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MCCUSKER ET AL.
1
The Climate Response to Stratospheric Sulfate Injections and Implications
for Addressing Climate Emergencies
KELLY E. MCCUSKER, DAVID S. BATTISTI, AND CECILIA M. BITZ
University of Washington, Seattle, Washington
(Manuscript received 1 April 2011, in final form 28 October 2011)
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ABSTRACT
AU1
Stratospheric sulfate aerosol injection has been proposed to counteract anthropogenic greenhouse gas
warming and prevent regional climate emergencies. Global warming is projected to be largest in the polar
regions, where consequences to climate change could be emergent, but where the climate response to global
warming is also most uncertain. The Community Climate System Model, version 3, is used to evaluate simulations with enhanced CO2 and prescribed stratospheric sulfate to investigate the effects on regional climate
and further explore the sensitivity of these regions to ocean dynamics by running a suite of simulations with
and without ocean dynamics.
The authors find that, when global average warming is roughly canceled by aerosols, temperature changes
in the polar regions are still 20%–50% of the changes in a warmed world. Atmospheric circulation anomalies
are also not canceled, which affects the regional climate response. It is also found that agreement between
simulations with and without ocean dynamics is poorest in the high latitudes. The polar climate is determined
by processes that are highly parameterized in climate models. Thus, one should expect that the projected
climate response to geoengineering will be at least as uncertain in these regions as it is to increasing greenhouse gases. In the context of climate emergencies, such as melting arctic sea ice and polar ice sheets and
a food crisis due to a heated tropics, the authors find that, while it may be possible to avoid tropical climate
crises, preventing polar climate emergencies is not certain. A coordinated effort across modeling centers is
required to generate a more robust depiction of a geoengineered climate.
1. Introduction
a. Motivation
The Intergovernmental Panel on Climate Change
Fourth Assessment Report (IPCC AR4) projects global
and annual mean warming of 1.78 to 4.48C this century
under the A1B emissions scenario (Meehl et al. 2007).
Warming in the northern high latitudes is projected to
be 1.5–4.5 times the global mean values in global climate
models (Holland and Bitz 2003). Even under stabilized
emissions or cessation of emissions, the planet will continue to warm due to the gases that have already been
emitted (Matthews and Caldeira 2008; Ramanathan and
Feng 2008; Solomon et al. 2009), with a very real possibility that committed warming is equal to or greater
than ‘‘dangerous’’ levels (Ramanathan and Feng 2008)
Corresponding author address: Kelly McCusker, Department
of Atmospheric Sciences, University of Washington, Seattle, WA
98195.
E-mail: [email protected]
DOI: 10.1175/JCLI-D-11-00183.1
Ó 2012 American Meteorological Society
that some claim may have catastrophic consequences
(Hansen et al. 2007). Blackstock et al. (2009) define climate emergencies as ‘‘those circumstances where severe
consequences of climate change occur too rapidly to be
significantly averted by even immediate mitigation efforts.’’ Such emergencies would include, for example,
the loss of habitat for polar bears, displaced arctic ecosystems, thawing permafrost, rapid sea level rise due to
melting Greenland and West Antarctic ice sheets, or a
large reduction in crop production due to temperature
changes in the tropics.
Recent Arctic warming and record summer sea ice area
minimums have spurred expressions of concern for, and
investigations into, the fate of sea-ice-dependent polar
bears (Regehr et al. 2010), arctic ecosystems (Grebmeier
et al. 2006), permafrost (Lawrence et al. 2008), and the
way of life of local communities (Hinzman et al. 2005 and
references therein). Each of these systems depends on
either sea ice area (e.g., for hunting, resting, breeding;
Moore and Huntington 2008) or subfreezing surface temperatures over land (e.g., permafrost and Greenland).
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Maintaining surface temperatures and preserving sea ice
are thought to be necessary to avoid threatening such
systems.
On the other side of the globe, the world’s greatest ice
sheets are found in West Antarctica and the Antarctic
Peninsula, storing huge amounts of water that could potentially raise sea level by many meters. There is a very
real danger that land ice calving and/or melting could
accelerate and cause greater sea level rise than is anticipated from thermal expansion of seawater alone. The
catalysts for such events are higher air temperatures,
and more importantly, warmer ocean waters sloshing
up against the ice sheet outlets that melt outlet glaciers
and ice shelves and could destabilize grounding lines
(Oppenheimer 1998; Schoof 2007). The most in-peril ice
sheets flow into the Ross and Weddell Seas (Oppenheimer
1998) in West Antarctica, which has been shown to be
warming currently (Steig et al. 2009) and losing ice mass
(Chen et al. 2009; Velicogna 2009).
Whereas the greatest warming is projected to occur
in the polar regions, the tropics show relatively modest
temperature changes under increasing CO2. Nonetheless, ecosystems in the tropics may be among those most
affected by a changing climate: small amplitude climate
variability in the tropics, combined with a tightly coupled ocean–atmosphere system, means that even small
climate changes can have important consequences for
living systems, whose evolution was built on a narrow
range of temperature. Organisms accustomed to stable
climatic conditions have lower physiological thresholds
and, thus, are put under more stress for a given warming than those from more climatically variable regions
(Tewksbury et al. (2008) and references therein). Specifically, crops grown in the tropics, providing livelihood
and sustenance to billions of people, abide by similar
laws, so even small climate changes in the tropics can
be detrimental. Global warming will cause temperature
and precipitation to surpass optimal growing conditions,
adversely affecting ecosystems, agriculture, and food
security for billions (Battisti and Naylor 2009).
Should climate evolve as models predict, and severe
consequences emerge, swift action may become necessary. However, even if all anthropogenic emissions of
greenhouse gases ceased, the planet would continue to
warm, possibly by a significant amount (Armour and Roe
2011; Hare and Meinshausen 2006). Thus the only solution in such a scenario would be to counteract the rise
in temperature by some other means. Accordingly, the
domain of feasible solutions to the global warming problem has expanded from adaptation and mitigation by
greenhouse gas emissions reductions to include geoengineering: the deliberate modification of the earth’s
radiative budget so as to stop the climate change due
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to increasing anthropogenic greenhouse gases. Geoengineering has evolved from a topic of intermittent
discourse between scholars (via publications), to news
media and the blogosphere. Recently, major world governments and important scientific societies—such as the
Royal Society of the United Kingdom (Shepherd and
Rayner 2009), the American Meteorological Society
(American Meteorological Society 2009), and the U.S.
Government Accountability Office—have made formal
statements and issued reports on the topic. Exploration of implementation and deployment technologies,
in some form or another, is currently being undertaken
(Blackstock et al. 2009), even to the point of pending
patent applications (Intellectual Ventures Laboratory,
Stratoshield White Paper, available online at http://
intellectualventureslab.com/wp-content/uploads/2009/10/
Stratoshield-white-paper-300dpi.pdf). Our current understanding of the effect geoengineering will have on the
climate system, especially on a regional scale, is not sufficient to rule out unfavorable consequences (Robock
et al. 2010), however, and calls for more research have
been made (American Meteorological Society 2009;
Hegerl and Solomon 2009; Keith et al. 2010).
Numerous schemes have been proposed to alleviate
the warming due to anthropogenic emissions of greenhouse gases. These schemes fall into two groups: those
that alter the sources and sinks of carbon in the earth
system, and those that alter the albedo of the planet.
Several implementation schemes have been proposed
to accomplish albedo management. These proposals include placing reflective mirrors in space, seeding clouds
to make them brighter, and injection of sulfate aerosols
or its precursors into the stratosphere. Each of these
schemes would work by allowing less shortwave energy
to reach the surface of the planet, thereby reducing
surface temperature. It is unknown which of these ideas
may be most effective in alleviating temperature rise.
However, sulfate injection is the leading contender because it is inexpensive to implement, uses existing technology (Robock et al. 2009), and it would be quick to
affect surface temperatures if commenced, as well as
quick to terminate (Matthews and Caldeira 2007; Robock
et al. 2008; Brovkin et al. 2009). These factors place
stratospheric sulfate injections on top of the list of relatively realistic solutions that could be deployed in the
near future and is the guiding reason for our choice to
simulate these injections in a GCM and study its effects
on the model’s regional climate.
b. Background
Many of the geoengineering modeling studies to date
evaluate the impact of aerosols or sunshade technology
by uniformly reducing the solar constant in a model.
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Govindasamy and Caldeira (2000) and Govindasamy
et al. (2003) showed in their studies that uniformly reducing the solar constant in an atmosphere general
circulation model (AGCM) coupled to a slab ocean
sufficiently canceled the global mean warming due to
doubling and quadrupling CO2, respectively. They found
a greater cooling in the tropics compared to the poles
and a reduction in seasonal amplitude of temperature.
In the context of the polar emergencies discussed in
the section 1a, these studies suffer from the exclusion
of ocean and sea ice dynamics. Lunt et al. (2008) performed a similar study, but with an AGCM coupled to
a full ocean GCM, albeit at low resolution, wherein they
quadrupled CO2 and reduced the solar constant. They
noted a reduced pole-to-equator temperature gradient,
reduced temperature seasonality, and a reduced intensity of hydrologic cycle compared with a preindustrial
control. Matthews and Caldeira (2007) investigated the
transient response of the climate to insolation reduction
in a model with an interactive carbon cycle, made up of
an energy–moisture balance atmosphere, dynamic ocean,
and dynamic–thermodynamic sea ice. Their study emphasized the fast response time of the climate to turning
on and off geoengineering. All of the above studies are
limited for the purpose of understanding the effects of
stratospheric aerosols by use of a forcing (reduced solar
constant) that has a very different spatial structure than
would be realized by stratospheric aerosol injections,
which we will show also has important implications for
the response of the climate system to the net forcing
of increased carbon dioxide plus aerosols.
The combination of imposed forcings is not necessarily, a priori, expected to result in stabilized climate
on a regional scale for three reasons. First, stratospheric
sulfate forcing, such as is prescribed in our experiments,
does not have the same properties as a forcing from
increased carbon dioxide because the former primarily
acts on shortwave radiation and the latter primarily on
longwave radiation. Thus, due to lack of sunlight, the
efficacy of sulfate aerosols in the polar regions may be
diminished. Second, studies have shown that modifications to shortwave versus longwave radiation affect
temperature and precipitation differently (Allen and
Ingram 2002; Bala et al. 2008). A perfect stabilization
of surface temperature by solar radiation management
necessarily excludes perfect stabilization of precipitation (both regionally and in the global average) because
of the differing energetic properties of the forcing agents.
Third, the spatial distribution of stratospheric sulfate
aerosol versus carbon dioxide is not identical, with carbon
dioxide being well mixed in the troposphere and a sulfate
layer limited to the lower stratosphere. The latter effect,
we will show, has profound implications for the response
3
of the climate—especially for the effectiveness of geoengineering to avoid the two polar emergencies that we
consider here. Even if the negating effect of a sulfate layer
was perfect, there is also some question as to just how
feasible it is to tune to the correct amount of sulfate in the
real world, where time scales of adjustment are spatially
varying and large natural variability will obscure the response of the earth system to changes in forcing.
Modeling studies of the response of the climate system to geoengineering have become more realistic by
including simulation of aerosol injection into the stratosphere (Rasch et al. 2008a; Robock et al. 2008; Jones
et al. 2010). Rasch et al. 2008a simulated geoengineering
by injecting sulfur dioxide into the stratosphere, in such
a manner that the quantity of particles would provide
enough (globally averaged) negative radiative forcing
to counteract the positive radiative forcing due to a
doubling of carbon dioxide. They found that the size
distribution of the aerosols affected the efficacy of the
cooling—specifically smaller sizes were more effective.
Robock et al. (2008) were the first to utilize a climate
model with a dynamical ocean and sea ice to investigate
the transient response to geoengineering with stratospheric sulfur injections. They simulated both tropical
and arctic instantaneous injections and found that the
effect of arctic injections, whose purpose was to recover
sea ice extent, was not restricted to the Arctic region but
extended south to 308N. Their model results also displayed a weakening of the Asian and Africa summer
monsoons. Though both Rasch et al. (2008a) and Robock
et al. (2008) use a more realistic forcing, in the context
of the climate emergencies in the polar regions, both of
these studies lack key ingredients: the model used by
Rasch et al. (2008a) did not include sea ice dynamics or
ocean dynamics, while the model employed by Robock
et al. (2008) is very insensitive to either greenhouse gases
or stratospheric aerosol (Jones et al. 2010) and the ocean
resolution used is extremely coarse. Ammann et al. (2010)
conduct fully coupled atmosphere–ocean model simulations that transiently counteract greenhouse warming (via
the IPCC A2 scenario) with either a solar reduction or
a stratospheric aerosol layer derived from the 1991 Mt.
Pinatubo eruption and focus on the regional effectiveness
of the combined radiative forcing. They find that the net
forcing induces enhanced atmospheric zonal circulation
anomalies, which contribute to residual Arctic warming.
The objective of this study is to investigate whether
the climate emergencies can be avoided through solar
management by injecting sulfate aerosols into the stratosphere. To our knowledge, this is one of the first attempts to maintain global mean surface temperature
transiently, in a fully coupled GCM with a more realistic
forcing (see also Ammann et al. 2010), and the only
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TABLE 1. Experimental details. The run names can be understood in the following way: ‘‘control’’ is the control run, ‘‘aero’’ has
a prescribed sulfate layer that is annually periodic in DISOM and ramped in OGCM, and ‘‘co2’’ carbon dioxide doubled in DISOM and
ramped in OGCM. DISOM is the Dynamic Sea Ice Slab Ocean Model and OGCM has both dynamic sea ice and ocean. Columns specify
the experiment name, type of ocean component, carbon dioxide concentration, and annual mean total atmospheric burden of sulfate, in
teragrams (1012g) of sulfur equivalent. The value of 1 3 CO2 is 355 ppm.
Expt
Ocean type
CO2 concentration
Sulfate aerosol burden(TgS)
control
aero
co2
geoco2
DISOM
1 3 CO2
1 3 CO2
2 3 CO2
2 3 CO2
0
8
0
8
control
aero
co2
geoco2
OGCM
1 3 CO2
1 3 CO2
1% ramp from 1990 concentration
1% ramp from 1990 concentration
attempt to illuminate the role of ocean dynamics in determining the regional response. We do this by ramping up carbon dioxide concentration concurrently with
stratospheric aerosol concentration, in our case with a
prescribed sulfate burden, in a climate model with full
ocean dynamics. We then compare these experiments to
experiments performed without full ocean dynamics to
illuminate some of the uncertainties in projecting a potential future geoengineered world that are particularly
relevant to the climate emergencies. We will focus on
three main regions of concern: the Arctic circle, the
West Antarctic ice sheet region (including the Antarctic
Peninsula), and the tropics. The paper is organized as
follows. Section 2 describes the global climate model and
experiment design. Results are described in section 3.
We first discuss the transient simulations and broad
global results. We further consider the vertical structure
of the atmospheric response to stratospheric aerosols
and to increased carbon dioxide and then discuss regional surface climate response to these forcings in the
context of the three climate emergencies. We expound
upon the uncertainties and discuss the broader implications of our results in section 4. Conclusions are presented in section 5.
2. Model and simulations
We perform our experiments using the National Center for Atmospheric Research (NCAR) Community Climate System Model, version 3 (CCSM3) (Collins et al.
2006), which has components for atmosphere, ocean,
land, and sea ice. We run each simulation with T42 resolution (approximately 2.88) in the atmosphere and a nominal 18 resolution in the ocean. The atmosphere has 26
vertical levels while the ocean has 40 vertical levels.
We run a suite of simulations with the atmosphere
component of the CCSM3 coupled to either a slab ocean
0
Linear ramp from 0
0
Linear ramp from 0
or to the full ocean general circulation model (OGCM)
of CCSM3 to determine the effect of ocean dynamics on
the climatic response to geoengineering (see Table 1).
The slab ocean utilized is a modified version of the more
common slab ocean model with motionless sea ice.
Our version has the complete CCSM3 thermodynamic–
dynamic sea ice model, and we refer to it as the Dynamic
Sea Ice Slab Ocean Model (DISOM). This model was
introduced and used by Holland et al. (2006) and Bitz
et al. (2006). The full atmosphere–ocean general circulation model configuration of CCSM3 is called OGCM
in this paper. The ocean heat flux convergence (OHFC)
prescribed in the DISOM simulations is derived from
the surface flux and ocean heat storage climatology of
the full OGCM from a 1990s CCSM3 control so that the
mean state of the DISOM and OGCM control simulations are the same. We use the ocean component (i.e.,
DISOM and OGCM) to differentiate between the model
configurations when referring to simulations (see Table 1).
We conduct experiments with various carbon dioxide
concentrations, some in combination with geoengineering. We do this using both configurations of CCSM3,
(i) DISOM (equilibrium) and (ii) OGCM (transient),
so that we conduct a total of eight simulations, listed in
Table 1. For each model configuration we have a control run (annually periodic external forcing from 1990s
levels, with CO2 5 355 ppm and other greenhouse gases
set to 1990 levels), an increased CO2 run (co2), a stratospheric sulfate-only run (aero), and a ‘‘net’’ run that has
both increased CO2 and a sulfate layer (geoco2).
The forcings are applied instantaneously in the DISOM
experiments, and then we run the model to equilibrium
(a minimum of 40 years). We analyze the last 40 years
of the DISOM control and geoco2 runs and the last
20 years of the co2 and aero runs. The CCSM3 OGCM
control and co2 runs were obtained from NCAR (Collins
et al. 2006). Carbon dioxide concentration is ramped
T1
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FIG. 2. The annual-mean zonal mean sulfate concentration multiplied by layer thickness (kg m22) at 2 3 CO2. The corresponding
global total burden is 8 TgS.
FIG. 1. Time series of global-mean annual-mean (a) surface
temperature (8C), (b) precipitation (mm day21), and (c) TOA net
flux (positive downward, W m22) for OGCM experiments. Ramping of sulfate and/or carbon dioxide begins in year 10 in the figure.
The 1870s preindustrial control is included for reference (black,
dashed line). The boxes indicate the years of averaging for all
results, unless stated otherwise in the text, red and blue dashed:
70–89, and black and green dashed: 60–99.
F1
F2
at 1% per year from the 1990 level. We ramp the sulfate
burden linearly from zero so that the amount prescribed
at any given time provides a global average negative
radiative forcing that approximately equals the positive
radiative forcing of the carbon dioxide. In the case of
geoco2, we integrated the model until the carbon dioxide reached four times modern concentrations of CO2
and the sulfate burden reached 16 teragrams of sulfur
equivalent (TgS). Figure 1a depicts the years for which
means are computed for each transient simulation. We
have one additional ensemble member for the OGCM
geoco2 simulation that was not available during the
initial analysis and writing of the paper. The ensemble
member exhibits essentially the same global mean changes
(to within 1%) and spatial pattern of response as the
geoco2 analyzed here, providing greater robustness to
our conclusions.
The sulfate forcing, or imposed ‘‘geoengineered layer,’’
is a prescribed burden of sulfate (SO4) in the stratosphere
and has a monthly climatology repeating annually. The
annually and zonally averaged sulfate concentration
multiplied by layer thickness at the time of CO2 doubling is shown in Fig. 2, which corresponds to a total
Fig(s). 1,2 live 4/C
annual mean burden of 8 TgS. By prescribing the aerosol
distribution, we ignore a major additional source of uncertainty in our study: the chemistry of sulfate formation
and its transport. However, these processes were taken
into account in the generation of the SO4 climatology.
The SO4 climatology is taken from the results of a
model study by Rasch et al. (2008a), whereby they
continuously injected a prescribed size distribution of
SO2 (sulfur dioxide) into the stratosphere at an altitude of 25 km from 108N to 108S, where it was transported by winds and interacted chemically. They used
a prescribed size distribution with a dry mode radius,
standard deviation, and effective radius values of 0.05,
2.03, 0.17 mm, respectively, which is meant to simulate
a volcanic-like size distribution. Once in the stratosphere,
the SO2 oxidizes to form sulfate aerosol, which is transported and removed via wet and dry deposition.
In the study by Rasch et al., the volcanically sized
aerosol distribution did not fully cancel the warming
due to doubled carbon dioxide. Because we prescribe
the aerosol distribution, we have scaled up the sulfate
climatology by the same fraction at each latitude and
height in the atmosphere, to better cancel the equilibrium warming under the 2 3 CO2 scenario experiment
in DISOM. This results in an annual mean prescribed
burden of sulfur equivalent in our simulations of 8 TgS
(to counteract 2 3 CO2) compared with 5.9 TgS in Rasch
et al. (2008a). It has been shown that there may be some
limitation to the effectiveness of sulfate aerosols when
the microphysics of sulfur dioxide injection and sulfate
aerosol formation are taken into account, such that the
burden required to cancel a doubling of CO2, for instance, would be greater than what is estimated in our
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TABLE 2. Global annual-mean (top) values and (bottom) differences (D) from respective control runs. OGCM means are calculated for
the 40 years surrounding the time of CO2 doubling in the control and geoco2 runs. The OGCM co2 and aero cases are 20-yr means
surrounding the time of CO2 doubling. Surface temperature (K), precipitation rate (mm day21), total Northern Hemisphere (NH) and
total Southern Hemisphere (SH) sea ice area (SIA) (1013 m2), total NH sea ice volume, total SH sea ice volume SIVol) (1013 m3).
Model
Expt
Surface temperature
DISOM
OGCM
control
control
288.1
288.0
DISOM
OGCM
DISOM
OGCM
DISOM
OGCM
geoco2
geoco2
aero
aero
co2
co2
Precipitation rate
2.80
2.79
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1.11, 1.23
1.12, 1.37
2.75, 2.07
2.74, 2.45
%D Precipitation
%D SIA NH, SH
%D SIVol NH, SH
0.09
0.08
23.36
21.27
2.70
1.38
21.47
21.43
28.54
23.70
5.21
2.15
23.89, 20.10
22.17, 22.21
25.5, 135
15.4, 20.19
228.8, 245.3
225.8, 210.3
29.59, 1.98
23.12, 23.09
128.6, 123.6
65.9, 3.03
252.0, 222.4
257.3, 215.7
3. Results
Our suite of experiments shows the extent to which
global and annual mean warming from rising CO2 can be
offset by placing sulfate aerosols with particular optical
properties and spatial distribution in the stratosphere.
We first show global-mean, annual-mean results and
annual mean spatial maps as a baseline. We then turn
our focus to specific regions, namely, the Arctic, West
Antarctica and the Antarctic Peninsula, and the tropics,
to examine the results in the context of climate emergencies.
a. Global
F3
SIVol NH, SH
DSurface temp
study (Heckendorn et al. 2009), and some have suggested that directly injecting sulfuric acid vapor may
improve the mass to radiative forcing ratio Pierce et al.
(2010). These complications are ignored here, as we
prescribe sulfate aerosols with a particular distribution
and specified optical properties.
T2
SIA NH, SH
Table 2 lists globally, annual averaged values of temperature, precipitation, and sea ice area and volume
from the set of simulations. It is no surprise that the
equilibrium temperature change of the DISOM geoco2
case relative to the DISOM control is near zero (Table 2)
because we adjusted the concentration of aerosols specified in the DISOM geoco2 run through several iterations.
It is more remarkable that the transient warming in the
OGCM forced by ramping CO2 at the rate of 1% per
year, shown in Fig. 1a, can be effectively canceled up to
about the 70th year after forcing commencement (the
time of CO2 doubling) by linearly ramping sulfate aerosol
concentration in the stratosphere.
For reference, Figs. 3a and 3b show the spatial maps of
annual average surface temperature change between the
DISOM and OGCM co2 and control simulations and
Figs. 3c and 3d show the companion maps for geoco2.
In the annual mean, the presence of a sulfate aerosol
layer is able to cancel surface temperature rises due to
increased CO2 nearly everywhere but the Arctic, which
will be discussed in more detail in the Arctic subsection.
The OGCM exhibits a slight global mean warming at
the end of the analysis period (see Fig. 1a). Hence, we
also compute the linear temperature trend spatially for
the period before the global mean temperature diverges
from the control (years 11–80). Figure 4 shows the magnitude of the temperature change extrapolated to year
80 (the midpoint of the analysis period), computed using
the linear trend of annual mean global mean surface
temperature for years 11–80. The pattern that emerges
matches that seen in Fig. 3d, indicating that the spatial
pattern of response is fundamental to the combination
of increasing CO2 and increasing sulfate layer burden,
and is not influenced by the existence of a residual global
mean warming (0.088C in the 40-yr average).
When aerosol concentrations are designed to cancel
global warming, they do not also cancel global mean
precipitation changes (Bala et al. 2008; Robock et al.
2008; Ricke et al. 2010). Indeed, placing sulfate aerosols
in the stratosphere reduces precipitation more than the
precipitation increases on average from raising CO2.
Thus, the globally averaged precipitation rate declines
by between 1% and 2% at the time of CO2 doubling for
both transient and equilibrium geoco2 cases relative to
their controls (see Table 2). Figure 1b shows the change
in precipitation with time for the OGCM simulations.
Although the globally averaged surface temperature
stays nearly constant for the first 80 years of the geoco2
experiment, precipitation slowly declines with increasing sulfate burden. This result supports the theory put
forth in Allen and Ingram (2002), whereby longwave
and shortwave radiation affect precipitation and temperature differently.
F4
AU2
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FIG. 3. Annual-mean surface temperature (8C) difference between (a),(b) co2 and control and (c),(d) geoco2 and control for (a),(c)
DISOM (at equilibrium) and (b),(d) OGCM simulations (at time of CO2 doubling). In (a),(b), the dots indicate regions of significant
cooling at the 95% level based on a Student’s t test. In (c),(d), the dots indicate regions where there is significant warming or cooling at the
95% level.
Figure 1c displays the top of atmosphere (TOA) net
flux anomaly from the control simulation for the OGCM
co2, aero, and geoco2. The OGCM geoco2 has an annual mean TOA imbalance anomaly of 0.06 W m22
during our analysis period, while the DISOM geoco2
has an imbalance of less than 0.01 W m22, indicating
that the geoco2 simulations are well balanced and the
radiative forcing from the imposed sulfate layer successfully counterbalances that from increased CO2 during
the analysis period.
b. Vertical structure
The global average forcing at the top of the atmosphere in the geoco2 experiments is effectively zero until
the time of CO2 doubling, but there are important spatial differences, particularly in the vertical. We discuss
and show the DISOM results of the vertical and zonal
Fig(s). 3 live 4/C
mean temperature in this section as an example; however, the OGCM has a very similar vertical temperature
response to the combined CO2 and aerosol forcing.
Raising CO2 causes tropospheric warming and slight
to -no cooling in the lower stratosphere (Fig. 5a). The
sulfate aerosol concentration is at a maximum over the
tropics, where the original injection of sulfur dioxide in
the Rasch et al. (2008a) study was located, which causes
an increase in absorption of solar and infrared radiation
there compared to the control climate. By virtue of this
spatial distribution, the sulfate aerosol produces a local
warming maximum in the lower stratosphere over the
tropical region (Fig. 5b). The net result of doubled CO2
and a sulfate layer on zonal mean temperature is to leave
the troposphere much like the 1990s control (Fig. 5c). Yet
in the stratosphere, the cooling due to increased carbon
dioxide does little to abate the tropical stratosphere
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sulfates does not counteract the circulation anomalies
due to increased CO2. We will see in the following sections that these upper-level differences are indeed manifested at the surface as changes in climate (although the
surface wind response tends to be weaker in OGCM than
DISOM).
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c. The Arctic
FIG. 4. Linear trend in the OGCM geoco2 simulation for years
11–80 (before global mean surface temperature diverges from the
control) at year 80, the midpoint of our analysis period. This shows
the pattern and amplitude of warming expected at year 80, given
the trend computed in years when there is no global mean surface
temperature trend. The pattern and amplitude are similar to the
annual mean change in temperature between the OGCM geoco2
and control shown in Fig. 3d. Thus, the pattern of response is fundamental to the combination of the sulfate layer and increased
CO2, not to the weakening sulfate effect at the end of the OGCM
geoco2 analysis period.
sulfate-driven warming. These nonneglible changes in
the vertical structure of temperature in the atmosphere
cause noticeable differences to the zonal-mean wind field.
Figures 5d–f display the vertical structure changes
in zonal-mean zonal wind in the DISOM perturbation
experiments. In the annual mean, enhanced CO2 forces
the Southern Hemisphere (SH) polar stratospheric vortex to shift equatorward, while the Northern Hemisphere
(NH) polar stratospheric vortex displays a broad, weak
enhancement in the upper atmosphere (Fig. 5d). The
subtropical tropospheric jets and zonal-mean surface
winds change little due to doubled CO2. In contrast,
forcing by sulfate aerosols alone causes a clear poleward
shift of the stratospheric and tropospheric polar vortex
in the SH and a strengthening of the stratospheric polar
vortex in the NH (Fig. 5e). The net result of the combined forcings in geoco2 looks similar to that of the sum
of the two separate forcings. In the NH the polar vortex
is strengthened even more in geoco2 than in co2 (an
increase in mean zonal wind of about 30% at the peak
location compared to 20% in co2), and this strengthening is especially apparent in December–February (DJF)
(not shown) when the strengthening also extends down
to the surface. In the SH the net result is an equatorward
shift of the stratospheric polar vortex and a poleward
shift in the jet in the troposphere. The zonal mean temperature and wind response patterns look very similar in
the OGCM for the geoco2 case. Thus, the addition of
Fig(s). 4 live 4/C
Figures 6a and 6b display the change in the summer,
June–August (JJA), surface temperature in the geoco2
simulations as compared to the control integrations for
the DISOM and the OGCM, respectively. Also noted
on the figure is the location of the sea ice edge, defined as
the region within which there is a 15% or greater sea ice
concentration, for geoco2 (dashed) and control (solid).
As expected, summer temperatures over sea ice remain
unchanged, as ice keeps the air temperature at about 08C
in summer. However, northern land surfaces are overcooled in both models, at some locations by over 18C,
with the notable exception of warming in Greenland.
Although the sea ice edge in the Arctic is nearly unchanged, the NH sea ice volume is reduced by 10.0%
and 2.8% for DISOM and OGCM, respectively, with the
greatest thinning of sea ice found around Greenland in
the DISOM and East Siberian Sea in the OGCM (not
shown, but the pattern is very similar to DJF with slightly
greater magnitude changes–see Figs. 6e–f). Additionally,
there are reductions in sea ice concentration (not shown)
of up to 10% in the marginal ice zones, especially in
DISOM.
Figures 6c and 6d show the change in surface temperature in Northern Hemisphere winter (December–
February). Contrary to JJA, the DJF surface temperature
response in geoco2 displays a broad residual warming
throughout much of the Arctic. This is expected, since
the impact of aerosols is diminished in the polar night
and the warming due to increased CO2 is not compensated. This warming results in increased water vapor and
an increase in surface evaporation throughout the high
latitudes that further contributes to the surface warming by an enhanced greenhouse effect. However, there
is spatial structure to these temperature changes that
cannot be due to the well-mixed CO2, as both model
configurations exhibit enhanced regional warming of
28–48C over northern Eurasia. Although this is cooler
than a 2 3 CO2 world, temperature differences are still
up to 50% of the warming expected from 2 3 CO2 here,
with the broadest regional warming occurring in the
DISOM model (cf. Figs. 3a,b with Figs. 6c,d). The sea ice
in DJF is decreased in both area (along sea ice margins,
not shown) and volume of 3.5% and 8.8% respectively in
the DISOM and 2.1% and 2.8% in the OGCM, although
the sea ice extent is nearly unchanged. Figures 6e–f
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FIG. 5. Zonal-mean, annual-mean (a)–(c) temperature and (d)–(f) wind in the DISOM simulations. Contours are the control; in colors
are differences between the perturbed experiment and the control experiment. The thick black line indicates the zero line; dashed is
negative temperature or easterly wind anomaly and solid is positive temperature or westerly wind anomaly. The contour interval is 1.08C
or 8 m s21.
F7
display the associated pattern of sea ice thickness change,
exhibiting more widespread and greater thinning in the
DISOM than OGCM.
Compared with the control climate, both the DISOM
and the OGCM geoco2 simulations have enhanced winter west to southwesterly winds at 950 mb over northern
Eurasia and the northern Atlantic Ocean and Nordic
seas, with the OGCM 950-mb wind enhancement mostly
limited to the Nordic seas (Fig. 7). These increased winds
are the near-surface expression of the upper-level polar
vortex, which is enhanced in winter as described earlier. The 950-mb zonal wind changes are statistically
significant everywhere: the magnitude is about 1 m s21
or greater. The circulation anomalies enhance the advection of climatologically warmer and moister marine
and lower-latitude air into northern Europe and Asia
(v9 $T and v9 $q) and help to explain the robust pattern of surface warming over Eurasia. The anomalous
moisture transport increases downwelling longwave at
the surface directly and also by providing additional
water vapor for latent heating of the atmosphere. In the
OGCM, Eurasia is also warmed by increased sensible
Fig(s). 5 live 4/C
heat advection associated with the reduction in the sea
ice thickness and in the area covered by sea ice to the
north of Eurasia.
The pattern of atmospheric circulation change and
surface warming is familiar as the postvolcanic eruption
winter response in which the climate exhibits a positive
Arctic Oscillation (AO) phase due to a strengthened polar vortex (Robock 2000; Stenchikov et al. 2002; Shindell
et al. 2004). However, the pattern cannot uniquely be
attributed to the stratospheric aerosols in this case. The
sulfate alone induces strengthened westerlies at the surface, most strongly only over northern Eurasia in the aero
simulations (not shown), which implies that much of the
AO response pattern elsewhere is due to increased CO2,
as the co2 case also exhibits surface circulation changes
of the same sign as a positive AO (not shown). This
suggests that the sulfate layer does not counteract CO2induced circulation changes; rather it nudges the circulation in the high northern latitudes in the same way as
does increasing CO2. This anomalous circulation plays
a dominant role in structuring the pattern of temperature response in northern Eurasia through temperature
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FIG. 6. The difference in (a),(b) mean JJA and (c),(d) mean DJF surface temperature (8C) and (e),(f) DJF sea ice
thickness (m) between the geoco2 runs and their corresponding control runs. Contours are 15% sea ice concentration
in control (solid) and geoco2 run (dashed). Results are for (left) DISOM and (right) OGCM. Dots indicate regions of
significant warming or cooling at the 95% level using a Student’s t test.
F8
and moisture advection, which, in addition to local feedbacks, further amplifies downwelling longwave radiation in the region. The anomalous winds also exert a
wind stress on the ocean that affects ocean circulation in
the OGCM geoco2 simulation (see Fig. 8b). Note that the
zonal wind stress changes exhibited in Fig. 8 are comparable between the co2 (8a) and geoco2 (8b) simulations.
The spatial extent of polar residual winter surface
warming in the OGCM simulation is much smaller than
in the DISOM. In particular, the geoco2 OGCM exhibits cooler SSTs near the United Kingdom and south of
Greenland (where the DISOM does not), near regions of
Fig(s). 6 live 4/C
deep-water formation. This cooling in the North Atlantic
in the OGCM (which exists in the annual mean as well)
can only be due to changes in ocean heat transport—the
only difference between the two model configurations.
In Fig. 9, we show annual-mean, Atlantic and Arctic
Ocean zonal mean potential temperature and Atlantic
meridional overturning circulation (AMOC) anomalies in the NH for the geoco2 and co2 OGCM simulations.
The lower panel of Fig. 9b shows the Atlantic and Arctic
Ocean warming from increased CO2, which in some
places extends down below 2 km. The lower portion
of the lower panel in Fig. 9d shows the change in the
F9
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FIG. 7. Difference in mean DJF 950-mb winds between the geoco2 runs and their corresponding control runs for the (left) DISOM and
(right) OGCM models. In color, (a),(b) the climatological DJF surface temperature (8C) and (b),(c) the climatological DJF integrated
water vapor from the surface to 870 mb (kg m22).
F10
AMOC due to increased CO2, which is weakened everywhere. In the geoco2 simulation, the sulfate aerosols
have managed to cool the ocean everywhere (note the
reduced color scale), and particularly north of 508N
(upper portion of the upper panel of Fig 9b). The effect
of the combination of sulfates and increased CO2 on the
AMOC is slightly more complex. The AMOC is weakened poleward of 308N above 1 km in depth. The net
result of the ocean temperature and circulation changes
under geoengineering is a reduction in northward heat
transport in the NH, shown in Fig. 10 along with the co2
and aero anomalies. Thus, in the geoco2 scenario there
is less residual warming and thinning of sea ice in the
OGCM than in the DISOM model.
Fig(s). 7 live 4/C
In summary, arctic climate changes induced by increasing CO2 are not perfectly canceled by the injection
of stratospheric sulfate aerosols, especially in winter due
to the ineffectiveness of the sulfate aerosols when there
is no sunlight. There is a consistent warming signal over
northern Europe and Asia in geoco2 DJF in the DISOM
and OGCM that is enhanced by the intensified nearsurface westerlies over Europe and Asia. Changes in
the North Atlantic caused by the net forcings of sulfate
aerosol and increased CO2 are the same sign as, but
weaker than, changes occurring under increased CO2
alone. The residual surface winds give rise to circulation
and heat transport changes in the ocean. As a result, the
North Atlantic Ocean in the OGCM geoco2 experiment
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FIG. 8. Annual-mean OGCM zonal wind stress difference (dyn cm21) between (a) co2 and control and (b) geoco2 and control.
exhibits cooler surface temperatures and slightly thicker
sea ice regionally around Greenland and in the neighboring Arctic when compared to the DISOM, thus better
offsetting winter surface warming from increased CO2.
d. The Antarctic
F11
As in the Arctic, there is residual winter, June–August
(JJA), warming over Antarctica from the combined sulfate aerosol and CO2 forcing (Figs. 11a,b). Both DISOM
and the OGCM have residual warming on and around
the Antarctic Peninsula; however, the surface warming
is much more focused along the Antarctic Peninsula in
the DISOM, whereas it is widespread in the OGCM,
extending eastward past the Weddell Sea and along the
Antarctic shore to almost 1308E. It is these regions
where we focus our analysis. The DISOM and the
OGCM exhibit temperature changes of opposite sign
over East Antarctica. Atmospheric circulation differences in DISOM and OGCM, and associated ocean dynamical responses in OGCM, are responsible for these
differences in the surface temperature and sea ice responses between the DISOM and OGCM experiments.
The DISOM geoco2 case has strengthened westerlies
over the entire Southern Ocean (Fig. 11c), which is a
manifestation of the poleward shift in the subtropical jet
(Fig. 5f). The changes in atmospheric circulation lead
FIG. 9. Annual-mean OGCM Atlantic Ocean zonal mean potential temperature (8C) for (a) control and (b) perturbation experiment
differences. Atlantic Ocean meridional overturning circulation (Sv) for (c) control and (d) perturbation experiment differences. Note the
smaller color scale for geoco2 as compared to co2.
Fig(s). 8,9 live 4/C
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FIG. 10. Annual-mean Atlantic Ocean northward heat transport
in (a) the control and (b) change in annual mean OGCM Atlantic
Ocean northward heat transport (PW) for geoco2 (solid), co2
(dash-dot), and aero (dashed).
to advection of climatological warm air from the north
by the anomalous north–northwesterlies just west of the
Antarctic Peninsula in the DISOM, causing warming over
the Bellingshausen Sea and the peninsula. Additionally,
the wind stress causes sea ice to be transported away
from the east coast of the peninsula into the Weddell
Sea where it accumulates. This creates the pattern of
thinner ice adjacent to thicker ice in Fig. 11c. This thinner
ice allows more heat from the ocean to be conducted
through the ice to the air throughout the winter season.
Then the thinner ice is advected eastward by the mean
westerlies, generating the warming maximum along the
east side of the peninsula and extended warmth to the
northeast.
The responses in Antarctic winter to the combined
aerosol and CO2 forcing in the OGCM model is different from that in the DISOM model in two ways. First,
coupled feedbacks in the tropical Pacific cause changes
in the mean state that affect the tropospheric winds in
the SH via atmospheric teleconnections (also see section 3e). Second, the subsequent changes in wind stress
on the Southern Ocean force changes in the ocean circulation that greatly affect the upper-ocean temperature
and thus sea ice thickness. We break the Antarctic response into two main regions, depicted in Fig. 11d.
Region A lies west of the Antarctic Peninsula and encompasses the Bellingshausen and Amundsen Seas.
Region B lies east of the peninsula and is focused on
the Weddell Sea but extends east to roughly 708E.
13
The zonal surface wind and wind stress anomalies
that are evident in the DISOM geoco2 case are also
found in the OGCM, but they are overwhelmed by a
much larger eddy contribution in JJA. In particular, in
region A an anomalous high pressure region (seen as
anticyclonic wind stress circulation in Fig. 11d) exists in
austral winter. The geopotential height anomalies in the
OGCM extend in the vertical, are nearly equivalent
barotropic, and display a clear Rossby wave train signal
emanating from the tropics (not shown). A recent study
(Ding et al. 2011) has shown that such a high pressure
anomaly in the Bellingshausen and Amundsen Seas is
induced when tropical sea surface temperatures are prescribed in a GCM to have a warm anomaly in the central
Pacific Ocean in JJA. As is described in the next section, the OGCM geoco2 case does indeed display a warm
anomaly in the central Pacific, albeit small, but similar
in magnitude and positioning as in the Ding et al. (2011)
study (0.28–0.38C in the OGCM compared with 0.28–
0.48C in Ding et al.). Thus, the warm anomalies over
the western Antarctic seas and peninsula, and most
importantly the Ross ice shelf, are due to anomalous
atmospheric advection of mean temperature, and are consistent with tropical SST changes generating a Rossby
wave train that reaches the shores of Antarctica.
Region B exhibits more expansive thinning of sea
ice in the OGCM than in the DISOM everywhere from
708W to 608E (see Figs. 11c,d). Figure 12 displays OGCM
geoco2 ocean potential temperature differences at various depths in SH winter, along with a cross section of
the zonal average ocean temperature difference within
the longitude sector demarcated in the figure. North
of about 608S, there is warming at most depths as the
nearly vertical isotherms have shifted southward. South
of roughly 608S in the Weddell Sea there is residual
warming above about 200-m depth and cooling below,
suggesting that there is anomalous upward heat transport in that region. To diagnose the source of heat near
the surface that thins sea ice in the Weddell Sea, be it
the atmosphere or ocean, we next examine ocean temperature and heat transport anomalies and perform an
energy budget analysis.
We divide the Weddell Sea basin, defined by the
markings in Fig. 12b with a northern edge at 608S, into
upper-layer (0–200 m) and bottom-layer (200–5000 m)
boxes. We then compute the vertical energy flux between the layers as a residual after taking into account
the top surface fluxes, the horizontal fluxes, and the temperature tendency in the upper layer. The upper layer
displays a positive temperature tendency, but a horizontal heat divergence and reduced top surface heat
flux into the layer. Thus, energy conservation demands
that there is an anomalous upward flux of heat into
F12
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FIG. 11. (a),(b) Mean JJA surface temperature change in geoco2 minus control (color shading) and sea ice extent defined as the 15%
concentration contour in control (solid) and geoco2 (dashed). Dots indicate regions of significant warming or cooling at the 95% level using
a Student’s t test. The sea ice extent in geoco2 and control are nearly equal. (c),(d) Mean JJA sea ice thickness change in geoco2 minus
control (color shading) with sea ice extent as in (a),(b). Vectors are differences in wind stress on the ocean: (left) DISOM and (right) OGCM.
the upper 0–200-m layer from below. We find a value of
1.3 W m22 upward flux into the upper layer from below.
The increase in upper-layer heat in the Weddell in the
geoco2 OGCM simulation is due to the poleward shift
of the Antarctic Circumpolar Current (ACC, not shown).
This poleward shift of the ACC in the geoco2 OGCM
simulation (not shown) appears as an enhancement in
zonal circulation at the northern edge of the Weddell
Sea, a greater Ekman transport northward, and hence
an increase in the divergence and upwelling of warmer
circumpolar deep waters south of the shifted ACC. South
of 558S in the Weddell Sea, isotherms are shallower in
the top 500 m and are generally displaced southward
in the geoco2 run, compared to the control run, at the
approximate latitude of the ACC in this region (see
Fig(s). 11 live 4/C
Fig. 12a). These changes are due to the increased zonal
wind stress on the ocean (shown in Fig. 8b) that occurs in
the annual mean (but is obscured in Fig. 11d because of
the large eddy component evident in JJA). Notably, the
geoco2 zonal wind stress anomalies around Antarctica,
shown in Fig. 8b, are similar in magnitude and pattern
to those in an increased CO2 world (Fig. 8a). Hence, as
in the Arctic, CO2-induced circulation changes are not
canceled by the inclusion of sulfate aerosols.
e. The tropics and subtropics
Figure 13 displays the tropical and subtropical (308S–
308N) seasonal surface temperature changes in the geoco2
simulations as compared to respective controls for
DISOM and OGCM. Temperature changes within this
F13
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FIG. 12. (a) Annual-mean OGCM potential temperature difference (8C) between geoco2 and control zonally
averaged in the sector demarcated and shown at various depths in (b). Contours are control (solid) and geoco2
(dashed) potential temperatures.
F14
region for all models and seasons are smaller than 18C
and usually less than 0.58C. In fact, much of the temperature change in the tropics under geoengineering,
especially on land, is not significantly different from the
control climate at the 95% level, as judged by a Student’s
t test. However, there is a sizable cooling over the equatorial Pacific in DJF and JJA in DISOM, which has an
effect on tropical precipitation.
Figure 14 shows the corresponding images for precipitation. As discussed earlier, in both DISOM and
OGCM there is a net drying due to the combined forcing
of aerosols and CO2, although there is interesting and
important spatial structure in the precipitation changes.
Over the oceans, the cooled regions are usually drier
regions (in general, the regions of statistically significant
precipitation change coincide with regions of statistically
significant temperature change). The control simulation
of CCSM3 features a double intertropical convergence
zone (ITCZ) over the Pacific Ocean: there are two
branches of high precipitation straddling the dry equator
(Collins et al. 2006) and both of our DISOM and OGCM
Fig(s). 12 live 4/C
control simulations display this behavior. Thus the changes
in precipitation in geoco2 in the DISOM translate to a
widening and strengthening of the dry tongue along the
equator. It is noteworthy that the sign of the precipitation
change along the equator in the central tropical Pacific
in the OGCM, although not statistically significant, is
opposite to DISOM, in JJA and particularly in DJF. The
local precipitation maximum here is consistent with the
local warming signal in the same location. This feature
also appears as a local warming maximum in the co2 runs
(not shown). As discussed in the previous section, this
feature in OGCM geoco2 is associated with the generation
of a teleconnection pattern in the southern Pacific Ocean
in JJA that produces the anomalous high pressure west
of the Antarctic Peninsula and thus the residual warming
seen in the Ross and Amundsen Sea regions due to anomalous atmospheric temperature advection (see Figs. 11b,d).
Ocean dynamics is clearly very important to the response
of the tropical pacific SST and precipitation over the ocean.
In terms of monsoonal precipitation, the DISOM and
OGCM model results in Fig. 14 clearly indicate an
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FIG. 13. The difference in (a),(b) mean JJA and (c),(d) mean
DJF surface temperature (8C) between the geoco2 runs and their
corresponding control runs: (a),(c) DISOM and (b),(d) OGCM.
Dots indicate regions of significant warming or cooling at the 95%
level using a Student’s t test.
increase in summer precipitation over western India.
East Asian coastal locations show a decrease in summer
precipitation, and there are hints of an African summer
monsoon reduction as well. However, the results are not
significant in these regions. This pattern of precipitation
response differs from Robock et al. (2008), which has
a weakened Southeast Asian monsoon and no increase
in Indian summer precipitation. In fact, across published
sulfate injection studies, there is little agreement in the
regional pattern of precipitation changes (Robock et al.
2008; Rasch et al. 2008b; Jones et al. 2010).
f. Beyond a 2 3 CO2 geoengineered world
We extended the OGCM geoco2 simulation past the
point of CO2 doubling to investigate whether the sulfate
layer is still able to counteract increased CO2 at higher
levels. Starting from doubled CO2 in the geoco2 run,
CO2 is increased 1% per year and the annual mean
sulfate burden is linearly increased until the concentration of CO2 is four times 1990 levels and the aerosol
burden is 16 TgS. The green line in Fig. 1a displays the
time series of global-mean, annual-mean surface temperature. Beyond the time of CO2 doubling (year 80
in Fig. 1), the effectiveness of the aerosol to offset the
warming decreases: the global average temperature increases by 0.38C by the time CO2 triples and by 0.68C by
the time it quadruples. Also, by 3 3 CO2, the precipitation has declined 1.8% and at 4 3 CO2, 2.2%. Not
surprisingly, the TOA radiative balance slowly becomes
more positive as well. Enhanced albedo feedbacks could
play a role, as there are slight reductions in surface
Fig(s). 13,14 live 4/C
VOLUME 00
FIG. 14. The difference in (a),(b) mean JJA and (c),(d) mean
DJF precipitation rate (mm day21) between the geoco2 runs and
their corresponding control runs for (a),(c) DISOM and (b),(d)
OGCM. Dots indicate regions of significant precipitation change at
the 95% level using a Student’s t test.
albedo and cloud cover in the geoco2 (not shown), but
these decreases do not display stronger rates of reduction after year 80. The increasing net TOA flux is thus
likely due to a slow saturation in sulfate scattering leading
to the appearance of nonlinearities in the burden to sulfate aerosol radiative forcing ratio at high sulfate burdens.
The studies of Matthews and Caldeira (2007), Robock
et al. (2008), and Jones et al. (2010) indicate that the
response time of the global mean surface temperature
to solar radiation management is relatively fast, and an
abrupt termination of aerosol injections would cause
a rapid rise in temperature back to where it would have
been had no geoengineering been implemented. We
performed one additional simulation where the sulfate
layer was abruptly shut off at the time of CO2 tripling, to
confirm previous work investigating the effects of a sudden termination of sulfate loading. The orange line in
Fig. 1a indicates that a sudden termination of geoengineering leads to a rapid temperature rise back to
what it would have been had such measures never been
performed. In this case, the rough estimate is that the
global mean temperature would rise 28C in a matter of
20 years. Hence, the rate of temperature rise is greatly
increased when compared to a scenario that never implements geoengineering.
4. Discussion
Our results show that climate change under stratospheric aerosols and increased carbon dioxide is smaller
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than under increased CO2 alone. However, maintaining
the modern climate is not possible: the combined forcings result in residual changes in the annual averaged
climate, in the seasonality of temperature and precipitation, and in regional patterns of atmospheric and oceanic circulation. Many of these residual differences result
because the aerosol layer is not able to counterbalance
the circulation anomalies induced under increased CO2.
Moreover, avoiding polar climate emergencies is not a
certainty, in part because CO2 forcing continues to operate throughout winter there, whereas the sulfate layer
is less effective.
We evaluated the role of ocean dynamics by comparing experiments with a climate model coupled to a
full ocean general circulation model and to a slab ocean.
Under the joint forcing of increased carbon dioxide and
an aerosol layer, surface temperature in northern Eurasia is cooled in summer, yet exhibits residual warming
in winter. The residual climate changes in the Arctic are
partially muted by ocean dynamical feedbacks that reduce the amount of poleward ocean heat transport into
the North Atlantic Ocean, limiting the thinning of sea
ice around Greenland and keeping SSTs cooler than when
the ocean dynamics is prescribed. The repercussions of
increased winter surface temperature in northern Eurasia
go beyond reduced sea ice. Arctic marine mammals in
general are not equipped to adapt swiftly to climate
changes (Moore and Huntington 2008) and thus their
well-being may be compromised by such a residual
warming.
With forcing by both increased CO2 and sulfate aerosols, Antarctica exhibits overcooling on the continent
in summer (not shown) and residual warming on and
around land surfaces in winter. Surface wind changes
drive changes in ocean circulation and act to amplify
the surface warming in winter around Antarctica. Anomalous upward ocean heat flux in the Weddell region results in slightly greater upper-ocean temperatures as
well. This residual upper-ocean and surface air warming
in and around the ice sheet exit regions does nothing to
allay the potential for the West Antarctic ice sheets to
become unstable due to increased melting, especially at
the base of the ice shelves (Oppenheimer 1998; Meehl
et al. 2007; Thoma et al. 2008; Jenkins et al. 2010). It is
this instability (tipping point) that causes concern about
rapid sea level rise (Notz 2009).
We find the highest effectiveness of geoengineering
for our climate emergencies is in the tropics. Except in
some places over the ocean, the temperature and precipitation differences under increased carbon dioxide
and a sulfate layer are small. This suggests that it may
be possible to avoid serious food security problems and
deleterious impacts on tropical organisms, so long as
17
the reduction in surface shortwave flux does not cause
adverse impacts on crop yields. Furthermore, models
tend to agree more in projections of future warming in
the tropical regions (see Fig. 15b), providing some confidence that these tropical projections of surface temperature would be robust to the choice of model used to
evaluate the impact of a geoengineering scenario. However, the inclusion of ocean dynamics is crucial, as even
the sign of the equatorial Pacific Ocean temperature
and precipitation response to increased CO2 and sulfate burden depends on ocean dynamics. This affects
the circulation response in the Southern Hemisphere
through teleconnections, which greatly affect the surface
air and ocean temperatures that bathe the ice shelves of
West Antarctica and the Antarctic Peninsula (Ding et al.
2011).
There is considerable uncertainty in other aspects of
stratospheric aerosol injections that are not related to
climate response per se: for instance, the creation of
the sulfate aerosol layer itself. Heckendorn et al. (2009)
have found that when using a 3D chemistry–climate
model with a 2D aerosol model to simulate sulfur dioxide injection into the stratosphere, the aerosol sizes
grow larger than expected. The result is an increase in
particle sedimentation and an ensuing nonlinear relationship between aerosol burden and injection rate,
resulting in even more aerosol being necessary to stabilize the global average temperature. Others have
found that increasing the sulfate burden in the stratosphere could delay the recovery of the ozone hole by
between 30 and 70 years due to the increased surface
area that the sulfate aerosol provides to catalyze ozone
destruction reactions (Tilmes et al. 2008).
Model uncertainties are amplified in the sensitive high
latitudes (Randall et al. 2007; DeWeaver 2007). These
uncertainties result in large differences in ocean and
ice variables among the IPCC AR4 models forced with
a business-as-usual greenhouse gas ramping scenario
(Meehl et al. 2007) and contribute to the larger spread in
polar climates among Coupled Model Intercomparison
Project 2 (CMIP2), which are due to wide ranges of
ocean heat transport, mean state of sea ice, and cloud
cover variables (Holland and Bitz 2003; Holland et al.
2001; Bitz 2008). Figure 15a displays the intermodel
average of the annual-average surface temperature difference between years 2080 and 2099 compared to years
1980–99 in the IPCC A1B emissions scenario among
AR4 models, while Fig. 15b displays the standard deviation in the ensemble average change in annual average surface temperature. The greatest differences in the
projected warming are found in the Arctic Ocean and
Southern Ocean, particularly in the Ross and Weddell
Seas under areas of current sea ice cover. These are
F15
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FIG. 15. The change in annual-average surface temperature simulated by the CMIP3 models used in the latest IPCC Assessment Report.
(a) The change over the twenty-first century (2080–99 mean minus 1980–99 mean), averaged across all the CMIP3 models, and (b) the
difference between the models, measured as the standard deviation of the simulated change from each model. Model output is from
forcing using the A1B emissions scenario.
the particular regions that may experience a climate
emergency.
Thus, putting our geoengineering simulations in the
context of the surface temperature spread in IPCC AR4
models, it is likely that the projected polar responses to
geoengineering will be highly sensitive to the choice of
the climate model. Accordingly, we endorse a call, put
forth by Kravitz et al. (2011), for modeling centers to
unite and execute a suite of coordinated, IPCC-style geoengineering simulations so as to sort out robust and
nonrobust responses to geoengineering. However, as
climate models fail to sample the long, low-probability
tail of very high future warming that is ubiquitous in estimates of climate sensitivity (Randall et al. 2007; Knutti
and Hegerl 2008; Roe and Baker 2007), they may also fail
to sample the full response to geoengineering.
5. Conclusions
This study fills a gap in research on stratospheric
aerosol injections by 1) imposing more realistic forcing
scenarios, 2) executing geoengineering simulations using
a standard resolution fully coupled atmosphere–ocean–
sea ice model with state-of-the-art sea ice physics, and
3) comparing the result with a mixed layer ocean model
with thermodynamic–dynamic sea ice to illuminate the
role of ocean dynamics. Importantly, we have also focused on regions that have the potential to experience
a climate emergency. We have executed a suite of experiments to simulate stratospheric aerosol injections
in a high CO2 world (assuming that the delivery system will deliver sulfate aerosols with specified optical
Fig(s). 15 live 4/C
properties). In general, and in keeping with previous
modeling work, the climatic effects of an aerosol layer
plus doubled carbon dioxide are smaller than in a world
with only doubled carbon dioxide. We have shown, however, that on seasonal time scales and regional spatial
scales, stratospheric sulfate does not necessarily cancel
all the effects of increased CO2, especially circulation.
In particular, there are still substantial climate changes
in the very regions where climate emergencies may drive
societies to geoengineer. Unfortunately, these are also
the regions that suffer the greatest uncertainty in the response to forcing, due to strong coupling between the
atmosphere, ocean, and sea ice and to deficiencies in
the parameterizations of unresolved physics in the models
(in particular, ocean mixing, sea ice rheology, atmospheric boundary layer processes, and clouds). Thus,
one cannot rule out the possibility of geoengineering
failing to avoid polar climate emergencies.
Countless other issues abound, both climatic and nonclimatic, including the ignorance of other consequences
of increased carbon dioxide, such as ocean acidification. The likelihood of a climate surprise occurring due
to geoengineering is high because research into geoengineering is still nascent, unintended consequences
are a certainty, and the uncertainties of geoengineering
are layered on top of those of global warming, compounding them. The question remains as to whether the
apparent global warming abatement geoengineering may
provide outweighs the (i) risk of foreseen consequences
being worse than predicted, (ii) risk of altogether unforeseen negative consequences, (iii) risk of failure in
international cooperation, (iv) risk of failure of the
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chosen geoengineering mechanisms, leading to rapid
temperature rise, and (v) risk of choosing winners and
losers in the climate battle. It is our opinion that it would
be imprudent to believe that the risk of unintended consequences is small enough to consider geoengineering a
solution at this time. More research is required, and a
coordinated modeling effort is a logical first step.
Acknowledgments. This research was funded by the
Tamaki Foundation and supported in part by the National Science Foundation through TeraGrid resources
provided by the Texas Advanced Computing Center
under Grant TG-ATM090059. We thank Philip J. Rasch
for providing data and for thoughtful discussions on experiment design and results, Kyle C. Armour for useful
comments on the manuscript, and Alan Robock and an
anonymous reviewer for suggestions that improved the
paper.
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